Magmatic history and evolution of continental lithosphere of the Sør Rondane Mountains, eastern Dronning Maud Land, East Antarctica

Precambrian Research 234 (2013) 63–84

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Magmatic history and evolution of continental lithosphere of the Sør Rondane Mountains, eastern Dronning Maud Land, East Antarctica Masaaki Owada a,∗ , Atsushi Kamei b , Kenji Horie c , Toshiaki Shimura d , Masaki Yuhara e , Kazuhiro Tsukada f , Yasuhito Osanai g , Sotaro Baba h a

Department of Earth Sciences, Yamaguchi University, Yamaguchi 753-8512, Japan Department of Geosciences, Shimane University, Matsue, Shimane 690-8504, Japan c National Institute of Polar Research, Tokyo 190-8518, Japan d Graduate School of Science and Technology, Niigata University, Niigata 950-2181, Japan e Department of Earth System Science, Fukuoka University, Fukuoka 814-0180, Japan f Nagoya University Museum, Nagoya 464-8601, Japan g Division of Earth Sciences, Faculty of Social and Cultural Studies, Kyushu University, Fukuoka 819-0395, Japan h Department of Natural Environment, University of the Ryukyus, Okinawa 903-0213, Japan b

a r t i c l e

i n f o

Article history: Received 29 February 2012 Received in revised form 8 January 2013 Accepted 19 February 2013 Available online 4 March 2013 Keywords: Pan-African suture Lithospheric evolution Geochemistry Zircon U–Pb dating Sør Rondane Mountains East Antarctica

a b s t r a c t The Sør Rondane Mountains, eastern Dronning Maud Land, East Antarctica, are situated within the PanAfrican suture zone, between West and East Gondwana, and the timing of collision event is regarded as the late Neoproterozoic to early Cambrian. In order to understand the tectonothermal history and evolution of the continental lithosphere, geochemical studies were conducted and zircon U–Pb SHRIMP dating was performed on intrusive rocks with basaltic compositions and associated metamorphic rocks. The metamorphosed tonalite complex exposed in the southwestern part of the mountains comprises tonalite associated with microgabbros, occurring as magmatic enclaves and later dikes that have intruded both the host and magmatic enclaves. Geochemically, the microgabbros are classified as low-Ti and high-Ti types, corresponding to the magmatic enclaves and dikes, respectively. Moreover, the geochemical features of the low-Ti microgabbro resemble those of an oceanic-arc tholeiite, whereas the high-Ti microgabbro has features of a back-arc basalt. The apparent zircon U–Pb ages show c. 990 Ma for the low-Ti microgabbro and c. 950 Ma for the high-Ti microgabbro. The high-grade metamorphic rocks situated in the northern part of tonalite complex were metamorphosed between 640 and 620 Ma, as constrained by zircon overgrowth rims of a migmatite leucosome (620 ± 2 Ma) and the previously reported age data. The timing of peak metamorphism corresponds to the early stage of the Pan-African suture event. Postdating the suturing event, unmetamorphosed minette dikes, dated 564 ± 2 Ma, intrude the tonalite complex and high-grade gneisses. On the basis of geochemical investigation, including Sr and Nd isotopic systematics, the microgabbros are considered to have originated from a depleted mantle source, whereas the minette magma is derived from an enriched mantle source. Consequently, the source mantles of the mafic magmas in the Sør Rondane Mountains have fundamentally changed from a depleted source in the early Neoproterozoic to a more enriched source in the late Neoproterozoic. Geochemical and isotopic evidence suggest that this compositional change of source mantle would reflect the interaction between the depleted mantle and the enriched crustal materials, such as subducted crustal rocks, caused by the Pan-African suture event. © 2013 Elsevier B.V. All rights reserved.

1. Introduction A multi-plate collision of various parts of East and West Gondwana produced the Pan-African suture during the Mesoproterozoic to Early Paleozoic. This suture zone stretches for more than 8000 km

∗ Corresponding author at: Yoshida 1677-1, Yamaguchi 753-8512, Japan. Tel.: +81 839335751. E-mail address: [email protected] (M. Owada). 0301-9268/$ – see front matter © 2013 Elsevier B.V. All rights reserved. http://dx.doi.org/10.1016/j.precamres.2013.02.007

and is recognized as the East-Africa-Antarctica Orogen (EAAO; Jacobs and Thomas, 2002, 2004) or the older East-Africa Orogeny (EAO; Stern, 1994), superimposed by the younger Kuunga Orogeny or Pinjara suture (Meert, 2003; Fitzsimons, 2003; Boger and Miller, 2004; Grantham et al., 2008). In this Orogenic belt, the older (c. 1000 Ma) basement rocks of Dronning Maud Land were differentially reworked during the Pan-African event. The Pan-African tectonothermal overprint was less intense in western Dronning Maud Land but strongly affected the Mesoproterozoic rocks in the region, extending from central to eastern Dronning Maud

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Fig. 1. Simplified geological map of the Sør Rondane Mountains (modified after Shiraishi et al., 1997). This map also shows the distribution of the microgabbros with their occurrence demarcated (magmatic enclave and dike).

Land. Moreover, the eastern end of Dronning Maud Land—the Lützow–Holm Complex—comprises juvenile Neoproterozoic crusts affected mainly by the Pan-African Orogeny with the Mesoproterozoic (c. 1000 Ma) tonalite, the Cape Hinode tonalite (Shiraishi et al., 1994). Voluminous intrusive rocks are exposed over a large area, extending from 2◦ E to 28◦ E. These intrusive rocks (600–500 Ma) are mainly undeformed, except in a few localized shear zones. The Sør Rondane Mountains (22 ◦ E–28◦ E) of eastern Dronning Maud Land are situated at the central part of the suture zone. This mountain range is underlain by early to late Neoproterozoic to early Cambrian metamorphic rocks and intrusive rocks with varied compositions (Fig. 1; Shiraishi et al., 1997). The peak metamorphic event occurred at around 640–600 Ma and has been regarded as the timing of the Pan-African suture event (Shiraishi et al., 2008). The igneous activities in the mountains represent pre- and post-collision stages. Since a mafic magma potentially preserves information on source mantle compositions, it can be used to analyze the evolution of the lithosphere from pre- to post-collision stages. Many researchers have reported metamorphic and magmatic ages from the Sør Rondane Mountains, as listed in Table 1. A metamorphosed tonalite complex (hereafter, tonalite complex), exposed in the southwestern part of the mountain range, represents the pre-collisional intrusive activity during the early Neoproterozoic (998–920 Ma; Table 1). Geochemical features of the tonalite complex show juvenile compositions (Takahashi et al., 1990; Tainosho et al., 1992). The bulk chemical compositions of some tonalites are similar to adakite, and are therefore probably derived from a subducted oceanic crust (Ikeda and Shiraishi, 1998). The tonalite complex contains a large amount of mafic rocks, lithologically referring to microgabbro. In general, the mafic intrusive rocks are important to obtain information on their source mantles; however, there are few detailed studies of this microgabbro. The high-K mafic dikes, minette and dolerite, related to the postcollision magma activities, intrude throughout the Sør Rondane Mountains. They discordantly cut the host rocks, and the timing of

magma activity ranges from 560 to 450 Ma (Table 1). These dikes show enrichment of large-ion lithophile element (LILE) and depletion of high field-strength element (HFSE) with negative Ti and Nb anomalies (Ikeda et al., 1995). The geochemical characteristics of the mafic dikes suggest that they are derived from an enriched source mantle (Ikeda et al., 1995; Owada et al., 2008, 2010). The microgabbros and high-K dikes represent the mafic intrusive rocks in the early Neoproterozoic and the late Neoproterozoic to early Cambrian, before and after the peak metamorphic event, respectively. Therefore, it is possible to analyze the chemical change of source mantles for the microgabbros and the mafic dikes beneath the Sør Rondane Mountains. The present study highlights the compositions of source mantle using geochemical and geochronological studies of mafic intrusive rocks during the Neoproterozoic Era. In addition, metamorphic rocks are dated to decide the precise age of peak metamorphic event. Finally, this paper discusses the magmatic history and evolution of continental lithosphere beneath the Sør Rondane Mountains. 2. Geology of the Sør Rondane Mountains The Sør Rondane Mountains consist of greenschist- to granulitefacies metamorphic rocks and various kinds of intrusive rocks (Shiraishi et al., 1991; Asami et al., 1992; Osanai et al., 1992; Tainosho et al., 1992). The main structural features of the metamorphic rocks are controlled by the east–west trend of the foliations and the fold axes (Toyoshima et al., 1995). The large east–west trending shear zone (Main Shear Zone) appears in the southwestern part of the mountains (Fig. 1; Shiraishi et al., 1991). The Sør Rondane Mountains was geologically divided into two groups, a layered gneiss group and a meta-tonalite (the tonalite complex) (Kojima and Shiraishi, 1986). The Main Shear Zone is situated between two groups (Fig. 1). Osanai et al. (1992) divided the metamorphic rocks into six units, from Unit I in the northeast to Unit VI in the southwest, in terms of constituent lighologies and geochemistry of mafic to intermediate gneisses of igneous protoliths. The tonalite complex corresponds to Unit VI.

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Table 1 Radiogenic ages for igneous and metamorphic rocks from the Sør Rondane Mountains. Lithology, occurrence

Method

Tonalite complex Tonalite, batholith Tonalite, batholith Tonalite, batholith Tonalite, batholith Tonalite, batholith Tonalite, batholith Tonalite, batholith Gabbro, stock Tonalite, dike Tonalite, batholith

U–Pb SHRIMP, Zircon U–Pb SHRIMP, Zircon U–Pb SHRIMP, Zircon U–Pb SHRIMP, Zircon U–Pb SHRIMP, Zircon U–Pb SHRIMP, Zircon U–Pb SHRIMP, Zircon U–Pb SHRIMP, Zircon U–Pb SHRIMP, Zircon Rb–Sr, whole-rock isochron

Layered gneiss Enderbitic gneiss Enderbitic gneiss Enderbitic gneiss Enderbitic gneiss Hbl-Bt gneiss Sil-Grt-Bt gneiss Grt-Bt gneiss Hb-Bt gneiss Bt gneiss Bt gneiss Grt-Bt gneiss Spl-Grt-Bt gneiss Sil-Grt-Bt gneiss Grt-Bt gneiss Grt-Bt gneiss Grt-Bt gneiss Sil-Grt-Bt gneiss Grt-Bt gneiss Grt-Bt gneiss Paragneisses

U–Pb SHRIMP, Zircon U–Pb SHRIMP, Zircon Sm–Nd, whole-rock isochron Rb–Sr, whole-rock isochron U–Pb SHRIMP, Zircon U–Pb SHRIMP, Zircon U–Pb SHRIMP, Zircon U–Pb SHRIMP, Zircon U–Pb SHRIMP, Zircon U–Pb SHRIMP, Zircon Th–U–Pb chemical age, monazite Th–U–Pb chemical age, monazite Th–U–Pb chemical age, monazite Th–U–Pb chemical age, monazite Th–U–Pb chemical age, monazite Th–U–Pb chemical age, monazite Th–U–Pb chemical age, monazite Th–U–Pb chemical age, monazite Th–U–Pb chemical age, monazite U–Pb SHRIMP, Zircon

Post-kinematic intrusive rocks Granite, stock Granite, stock Granite, dike Granite, stock Granite, stock Granite, stock Granite, stock Minette, dike Dolerite, dike Dolerite, dike Dolerite, dike Dolerite, dike Dolerite, dike

U–Pb SHRIMP, Zircon U–Pb SHRIMP, Zircon U–Pb SHRIMP, Zircon Rb–Sr, whole-rock isochron Sm–Nd, whole-rock isochron Rb–Sr, whoe-rock isochron Rb–Sr, whole-rock isochron K–Ar, biotite Ar–Ar, whole-rock K–Ar, whole-rock Ar–Ar, whole-rock K–Ar, whole-rock K–Ar, whole-rock

a b

Age (Ma) 995 ± 3 998 ± 3 995 ± 2 996 ± 3 945 ± 2 943 ± 3 920 ± 8 930 ± 10 772 ± 4 956 ± 39 951 ± 17a , b 602 ± 15b 961 ± 101a 978 ± 52a 1133 ± 12 637 ± 6 601 ± 6 605 ± 7 653 ± 11b 571 ± 5b 536 ± 4 542 ± 6 534 ± 6 534 ± 10 529 ± 14 543 ± 18 524∼550 539 ± 16 541 ± 15 1100–700 619 ± 7 564 ± 5 549 ± 13 528 ± 31a 519 ± 98a 525 ± 32 506 ± 43 563 ± 14 476 ± 23 467 ± 7 439 ± 13a 451 ± 12a 488 ± 18a

Interpretation

Reference

Intrusive age Intrusive age Intrusive age Intrusive age Intrusive age Intrusive age Intrusive age Intrusive age Intrusive age Intrusive age

Kamei et al. (2012) Kamei et al. (2012) Kamei et al. (2012) Kamei et al. (2012) Kamei et al. (2012) Kamei et al. (2012) Shiraishi et al. (2008) Kamei et al. (2012) Kamei et al. (2012) Takahashi et al. (1990)

Protolith age Metamorphic age Protolith age Protolith age Protolith age Metamorphic age Metamorphic age Metamorphic age Protolith age Metamorphic age Metamorphic age Metamorphic age Metamorphic age Metamorphic age Metamorphic age Metamorphic age Metamorphic age Metamorphic age Metamorphic age Inherited ages

Shiraishi et al. (2008) Shiraishi et al. (2008) Shiraishi and Kagami (1992) Shiraishi and Kagami (1992) Shiraishi et al. (2008) Shiraishi et al. (2008) Shiraishi et al. (2008) Shiraishi et al. (2008) Shiraishi et al. (2008) Shiraishi et al. (2008) Asami et al. (1996) Asami et al. (1996) Asami et al. (1996) Asami et al. (1997) Asami et al. (1997) Asami et al. (1997) Asami et al. (1997) Asami et al. (1997) Asami et al. (1997) Shiraishi et al. (2008)

Intrusive age Intrusive age Intrusive age Intrusive age Intrusive age Intrusive age Intrusive age Intrusive age Intrusive age Intrusive age Intrusive age Intrusive age Intrusive age

Li et al. (2006) Shiraishi et al. (2008) Shiraishi et al. (2008) Tainosho et al. (1992) Arakawa et al. (1994) Takahashi et al. (1990) Tainosho et al. (1992) Owada et al. (2010) Takigami and Funaki (1991) Takigami and Funaki (1991) Takigami and Funaki (1991) Takigami and Funaki (1991) Takigami and Funaki (1991)

The same sample determined with different methods. The same sample with different interpretation.

The tonalite complex is exposed in the southern part of the Main Shear Zone, whereas banded gneisses of felsic-to-intermediate compositions with minor amounts of pelitic and mafic rocks mainly crops out in the northern part of the Main Shear Zone (Fig. 1). The main litho-facies of the complex is tonalite. This tonalite complex was believed to be a single pluton; however, it is geochemically classified into tholeiitic and calc-alkaline series with different ages (Kamei et al., 2012; Table 1). According to Kamei et al. (2012), the timing of intrusion determined with zircon SHRIMP method is of 998–995 Ma for the tholeiitic tonalite and 945–920 Ma with minor 772 Ma for the calc-alkaline tonalite (Table 1). Both tonalites were derived from a juvenile environment related to subduction of oceanic crusts in terms of geochemical features. Osanai et al. (2013) proposed that the metamorphic rocks from the Sør Rondane Mountains can be divided into Northeastern terrane (NE-terrane) and Southwestern terrane (SW-terrane) in terms of their progressive metamorphism and P–T paths. The NE-terrane underwent granulite-facies metamorphism with minor amounts of amphibolite-facies retrogression and the metamorphic grade roughly increases toward the south (SW-terrane). On

the other hand, SW-terrane comprises granulite-, amphibolite-, and greenschist-facies metamorphic rocks. The metamorphic grade of SW-terrane increases toward NE-terrane from south to north. The P–T paths of granulites from both terranes are dissimilar, i.e. clockwise and counter-clockwise paths for NE- and SW-terranes, respectively. Subsequent to the metamorphic event, post-tectonic intrusive rocks intrude both terranes (e.g. Shiraishi et al., 1997; Li et al., 2003, 2005; Shiraishi et al., 2008). The boundary between NE- and SW-terranes would, therefore, be of a significant boundary, probably considered as a collision boundary (Osanai et al., in 2013). The post-tectonic intrusive rocks, granitic stocks and mafic dikes, intrude throughout the Sør Rondane Mountains. The granite stocks are massive and discordantly intrudes into the host gneisses and, locally, gives distinct thermal effects to the host gneisses (Asami et al., 1992). In the central part of the Sør Rondane Mountains, granitic suites are lithologically represented by biotite granite. The biotite granite is of evolved alkali–granite with high concentrations of K, Ba, and Sr (Takahashi et al., 1990; Tainosho et al., 1992; Li et al., 2001, 2005; Owada et al., 2006). The mafic dikes,

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Fig. 2. Occurrence of metamorphic and intrusive rocks from the southwestern part of Sør Rondane Mountains. (a) Felsic and mafic schist blocks in the host tonalite. (b) The tonalite intruding layered gneiss. Note that dark and light colored layers correspond with the host gneiss and the tonalite, respectively. (c) The tonalite complex consisting of the tonalite (white colored rock) and microgabbros (dark colored rock). Note: a microgabbro dike cutting microgabbro enclaves. (d) The migmatite with the pelitic gneiss. Note that the migmatite leucosome cutting pervasive foliations of the pelitic gneiss. (e) The minette dikes intruding a granite stocks with a shallow angle.

dolerite and lamprophyre (mostly minette), sporadically intrude the host gneisses (Ikeda et al., 1995). These dikes intrude in a within-plate tectonic setting, and the timing of magmatic activity begins after the suture event related to the formation of the Gondwana supercontinent (Ikeda et al., 1995; Owada et al., 2008, 2010). 3. Description of analyzed samples 3.1. Tonalite complex The southwestern part of the Sør Rondane Mountains is dominated by the tonalite complex, which consists mainly of tonalite accompanied by a microgabbro, and locally includes felsic and mafic schists as xenoliths with tabular shapes (Fig. 2a). The tonalite complex intrudes the layered gneiss, at places (Fig. 2b),

but it is mainly bounded by the gneisses at the Main Shear Zone (Fig. 1). The tonalite generally has pervasive foliations and is composed mainly of plagioclase, quartz, biotite, and hornblende with traces of K-feldspar, garnet, and epidote. Most of these minerals have been recrystallized and/or chemically modified during metamorphism; for example, epidote replaces or overlaps plagioclase. As described in the previous section, the tonalite geochemically belongs to the tholeiitic and calc-alkaline series, but both types are indistinguishable mineralogically. On the other hand, the degree of deformation is different; the tholeiitic tonalite locally shows strong foliations in contrast to the calc-alkaline tonalite (Kamei et al., 2012). The microgabbro is ubiquitous in the tonalite complex, occurring as the following two types: (1) magmatic enclaves (or syn-plutonic dikes) and (2) dikes clearly cutting the host tonalite

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Fig. 3. Photomicrographs of metamorphic and intrusive rocks. All photos are taken in the open nicol. (a) The microgabbro consisting recrystallized minerals. (b) The felsic schist showing strong deformation. (c) The migmatite leucosome containing kyanite. (d) The texture of minette shows that it is free from any deformational and recrystallized features. Hbl, hornblende; Bt, biotite; Epi, epidote; Chl, chlorite; Ky, kyanite; Pl, plagioclase; Qtz, quartz; Kfs, K-feldspar; and Ap, apatite.

(Fig. 2c). Both types of microgabbro show weak deformation and consist mainly of plagioclase, hornblende, epidote with traces of biotite, and quartz with few zircon grains (Fig. 3a). In addition, the microgabbro occurring as the dike locally contains titanite. As described later, bulk titanium content of the dikes are higher than those of the magmatic enclaves. Felsic and mafic schists are found in limited regions, mainly Walunumfjellet and the eastern part of Widerøefjellet. They occur as alternating layers and are included in the tonalite complex. The felsic schist consists of quartz, plagioclase, chlorite, and epidote with traces of biotite (Fig. 3b), whereas the mafic schist is composed of plagioclase, chlorite, epidote, and quartz. These mineral assemblages are similar to those of mylonitized parts of the tonalite (Shiraishi et al., 1992). The felsic schist includes zircon as an accessory mineral. 3.2. Kyanite-bearing migmatite Retrograde kyanite-bearing pelitic gneisses are exposed in the Koyubi-ridge of Brattnipene, the central part of the Sør Rondane Mountains Asami and Shiraishi (1987) and Asami et al. (1992). The migmatite in this study was collected from the Koyubi-ridge close to the location of the retrograde kyanite-bearing pelitic gneiss, described by Asami and Shiraishi (1987). The leucosome of garnetand kyanite-bearing migmatite locally cuts the pervasive foliations of the pelitic gneiss (Fig. 2d). This leucosome consists of quartz, plagioclase, K-feldspar, biotite, muscovite, kyanite, and garnet with minor amounts of zircon (Fig. 3c). The melanosome of migmatite is composed of garnet, biotite, sillimanite, kyanite, spinel, cordierite, quartz, plagioclase, and K-feldspar. These assemblages are similar to the pelitic gneiss described by Asami and Shiraishi (1987). On the basis of mineral texture, equilibrium mineral assemblage of garnet, sillimanite, biotite, K-feldspar, quartz,

and plagioclase is observed. Sillimanite occurs as an inclusion in garnet porphyroblasts and matrix, whereas kyanite appears only in the matrix, thereby suggesting isobaric cooling (Baba et al., 2012). 3.3. Minette Minette was collected from two localities, Lunckeryggen and Widerøefjellet, in the southwestern part of the Sør Rondane Mountains (Fig. 1). At both locations, the minette occurs as a dike, ranging from a few centimeters to meters in width (Fig. 2d). Generally, the dikes show dark color and have chilled margins. The minette from Lunckeryggen intruded both the tonalite and the Lunckeryggen granite stock, which is one of the post-collision granites in the Sør Rondane Mountains (Owada et al., 2006). The dikes strike N–S and gently dip to the west (20◦ –30◦ ). The minette was intruded by aplitic veins derived from the granite stock. Its intrusive age would therefore be similar to that of the granite stock (Owada et al., 2010). The minette apparently exhibits an igneous texture and consists mainly of green amphibole, greenish brown biotite, epidote, and feldspars. Apatite, zircon, and opaque minerals occur as accessory phases (Fig. 3d). The minette from Widerøefjellet described by Owada et al. (2010) intruded along normal faults in the tonalite complex in the southeastern end of Widerøefjellet. This minette possesses vesicles that are filled by calcite inside the minette. The dikes strike E–W and dip moderately to the south (40◦ –50◦ ). The minette bears phenocrysts of biotite, clinopyroxene, and feldspars. Biotite phenocrysts exhibit almost euhedral shapes and are fresh. Olivine phenocrysts are locally identified by their euhedral shape and are completely replaced by amphibole that probably formed during cooling. Opaque minerals, zircon, and apatite occur as accessory minerals.

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Table 2 Representative bulk chemical compositions. Sample

Locality

XRF (wt%) SiO2

TiO2

Al2 O3

Fe2 O3

MnO

Walnumfjellet Nils Larsenfjellet Nils Larsenfjellet Nils Larsenfjellet Lunckeryggen

49.18 54.15 52.93 48.71 49.06

0.56 0.48 0.40 0.54 0.63

17.94 15.23 15.21 14.60 14.59

12.25 10.24 7.17 10.24 9.97

0.22 0.27 0.18 0.28 0.22

Microgabbro (high-Ti) M08122101F M08122601B M08122602D M09010702 M09011301B

Nils Larsenfjellet Nils Larsenfjellet Nils Larsenfjellet Walnumfjellet Walnumfjellet

48.65 48.45 48.56 52.75 47.34

1.41 1.37 1.85 2.71 1.79

17.60 16.85 17.00 13.21 16.32

11.45 11.39 13.97 17.93 13.46

Minette M08121201B M08121201C M09010801 M09010802A

Widerøefjellet Widerøefjellet Lunckeryggen Lunckeryggen

47.87 47.18 53.62 49.23

1.05 1.02 1.34 1.62

13.12 13.02 13.47 13.17

Mafic schist M08122602B M09012401C M09012401J

Nils Larsenfjellet Widerøefjellet Widerøefjellet

45.38 52.72 48.68

0.42 0.50 0.38

Felsic schist M09012401G M09012401K1 M09012401L

Widerøefjellet Widerøefjellet Widerøefjellet

75.45 72.34 70.78

Migmatite & pelitic gneiss Brattnipene O90123101Ky1-2 O90123101Ky2 Brattnipene M08122801E Widerøefjellet M09011602F Walnumfjellet Walnumfjellet M09012301H Brattnipene O90123101A O90123101Ky1-1 Brattnipene

Microgabbro (low-Ti) M09010703A M08122301D M08122602A M08122602F M09011101

Sample

MgO

CaO

Na2 O

K2 O

P2 O5

Total

7.04 5.55 6.81 8.91 9.39

11.40 9.82 12.09 12.22 9.91

2.09 1.94 3.27 2.38 3.01

0.34 0.43 0.10 0.24 1.20

0.03 0.03 0.05 0.05 0.05

101.03 98.13 98.20 98.16 98.02

1.57 1.66 0.95 1.03 0.96

0.19 0.20 0.19 0.26 0.20

8.11 7.87 5.22 4.47 7.68

10.21 9.84 7.98 8.19 9.80

2.26 2.63 2.19 1.76 2.92

0.38 1.16 2.24 0.30 0.59

0.10 0.12 0.19 0.27 0.15

100.36 99.87 99.38 101.86 100.25

1.27 1.30 2.41 3.61 1.58

8.82 9.66 6.62 8.04

0.18 0.20 0.11 0.15

9.60 10.43 5.36 5.37

10.76 9.63 6.78 9.47

1.14 1.30 1.32 0.93

6.23 6.25 7.67 8.12

0.60 0.61 0.89 1.17

99.37 99.28 97.18 97.26

0.83 0.83 1.11 1.35

17.18 17.89 16.94

12.02 13.20 11.17

0.21 0.16 0.21

9.60 6.23 8.06

10.88 6.55 11.42

1.92 1.16 1.40

0.47 0.03 0.06

0.01 0.03 0.02

98.09 98.48 98.34

1.13 1.91 1.25

0.24 0.34 0.26

12.38 13.66 14.26

2.85 3.13 4.30

0.04 0.07 0.08

0.64 0.07 0.83

5.17 8.86 3.93

3.34 2.13 3.93

0.09 0.02 0.38

0.04 0.06 0.05

100.25 100.68 98.80

4.03 38.59 4.66

65.11 71.85 57.51 60.29 72.55 52.26 51.02

0.97 0.33 2.03 0.60 0.77 1.07 1.39

15.22 14.49 15.93 15.99 12.41 22.06 23.71

7.62 2.93 11.00 7.76 5.82 9.80 9.83

0.11 0.05 0.15 0.17 0.07 0.13 0.10

2.35 0.77 2.59 4.66 1.77 2.83 2.79

0.92 1.28 5.34 3.07 1.57 1.11 1.53

2.00 2.90 2.26 2.41 3.10 2.22 2.41

5.22 4.83 1.71 3.25 1.96 6.66 4.70

0.03 0.03 0.52 0.05 0.04 0.06 0.05

99.53 99.45 99.05 98.25 100.06 98.20 97.52

Hf

Nb

Ni

Rb

Sr

Ta

ICP-MS (ppm) Ba

Cr

Microgabbro (low-Ti) M09010703A M08122301D M08122602A M08122602F M09011101

46 91 32 68 183

50 30 310 560 460

0.5 1 0.5 2.1 1

0.6 1.6 <0.2 2.1 1.8

20 20 80 140 120

5 6 1 2 33

146 134 134 181 428

<0.01 0.08 0.02 0.17 0.08

0.26 1.23 0.81 4.58 0.57

Microgabbro (high-Ti) M08122101F M08122601B M08122602D M09010702 M09011301B

72 277 248 42 147

90 80 <20 <20 60

2.3 2.2 3.5 4.7 2.7

5.9 4.9 7.6 4.2 5.5

130 150 50 40 80

7 40 114 5 12

299 172 378 108 285

0.3 0.36 0.49 0.25 0.41

Minette M08121201B M08121201C M09010801 M09010802A

2520 4370 7400 8580

440 430 130 50

4.8 5.9 17.4 13.2

13.2 13.8 14 18.5

50 70 50 <20

147 236 224 264

660 631 2030 1790

Mafic schist M08122602B M09012401C M09012401J

56 17 21

320 40 70

0.4 0.5 0.2

0.7 0.4 0.6

110 30 40

12 2 4

35 25 165

<20 <20 <20

2.4 0.8 1.4

1 0.5 1.3

<20 <20 <20

Migmatite & pelitic gneiss 927 O90123101Ky1-2 921 O90123101Ky2 758 M08122801E 622 M09011602F 198 M09012301H

90 <20 <20 <20 50

12 4.9 11.6 3.6 7.3

25.1 7.3 23.3 4 9.2

30 <20 20 <20 20

Felsic schist M09012401G M09012401K1 M09012401L

FeO*/MgO

Th

U

V

Y

Zr

Rb/Sr

0.24 0.67 2.62 2.39 0.3

295 111 199 250 193

10.2 26.3 10.1 12.3 13.5

17 31 18 77 30

0.034 0.045 0.007 0.011 0.077

0.72 1.11 1.25 1.39 0.59

0.11 0.41 0.63 2.82 0.24

212 201 212 404 224

23.8 22.9 25.9 46.5 28.2

94 86 152 189 110

0.023 0.233 0.302 0.046 0.042

0.86 0.69 0.86 1.04

9.12 9.55 16.7 8.47

4.97 10.6 5.65 21

151 153 118 123

25.2 30.2 35.1 46.9

193 242 701 590

0.223 0.374 0.110 0.147

143 195 178

0.05 0.03 0.03

0.54 0.48 0.69

0.35 0.23 0.21

254 207 244

12.3 8.7 8.6

16 18 5

3 1 12

179 161 138

0.05 0.08 0.13

0.8 0.91 1.64

0.47 0.48 0.67

16 46 19

18.1 13.1 19.2

96 28 49

180 122 63 37 119

233 290 288 84 61

1.14 0.34 1.35 0.11 0.77

9.47 1.33 7.53 0.36 18.1

2.59 0.84 2.95 0.26 4.27

125 28 57 65 70

55.9 21.6 52.8 18.6 46.4

442 184 576 151 302

0.773 0.421 0.219 0.440 1.951

M. Owada et al. / Precambrian Research 234 (2013) 63–84

69

Table 2 (Continued ) Sample

ICP-MS (ppm) Ba

Cr

O90123101A O90123101Ky1-1

953 949

90 140

Sample

La

Ce

Hf 7.8 9.6

Pr

Nd

Nb

Ni

Rb

Sr

21.3 47.4

30 40

256 242

231 214

Gd

Tb

Sm

Eu

Microgabbro (low-Ti) M09010703A 4.20 M08122301D 3.30 M08122602A 2.80 8.27 M08122602F 12.9 M09011101

8.50 13.2 6.09 23.0 33.1

1.38 1.62 0.820 3.48 4.28

6.11 8.69 3.85 16.5 16.6

1.65 3.07 1.17 3.81 3.07

0.428 0.898 0.417 0.713 1.32

1.66 3.84 1.41 2.91 2.68

0.290 0.740 0.260 0.430 0.420

Microgabbro (high-Ti) 8.56 M08122101F M08122601B 9.09 16.9 M08122602D 5.15 M09010702 M09011301B 8.51

19.5 21.9 42.6 15.7 22.1

2.95 3.09 5.65 2.62 3.30

13.9 14.6 24.9 14.7 16.0

3.89 3.90 5.83 5.57 4.52

1.31 1.34 1.90 1.92 1.56

3.95 4.11 5.45 6.83 5.17

0.720 0.710 0.870 1.310 0.880

3.93 2.64 3.96 5.41

10.8 9.59 12.2 17.3

Minette M08121201B M08121201C M09010801 M09010802A

47.4 45.7 61.3 73.8

Mafic schist M08122602B M09012401C M09012401J

4.18 2.63 2.36

Felsic schist M09012401G M09012401K1 M09012401L

7.18 4.88 7.50

Migmatite & pelitic gneiss 30.7 O90123101Ky1-2 O90123101Ky2 22.7 M08122801E 58.2 M09011602F 13.7 M09012301H 39.0 O90123101A 66.8 O90123101Ky1-1 101

100 97.0 134 164 9.05 5.90 3.42

13.3 13.0 17.7 22.3

56.6 53.7 73.4 94.8

12.5 11.6 15.3 20.8

1.29 1.28 1.64 2.33

Ta 1.28 2.51 Dy

Th

U

20.5 38.5

5.84 8.55

V

Y

Zr

Rb/Sr

148 199

64.4 96.7

326 369

1.108 1.131

Ho

Er

Tm

Yb

Lu

Ce/Sm

Sm/Yb

1.78 4.67 1.69 2.36 2.42

0.360 1.06 0.380 0.470 0.510

1.08 3.12 1.13 1.30 1.44

0.178 0.511 0.182 0.199 0.214

1.20 3.30 1.23 1.23 1.44

0.184 0.549 0.195 0.201 0.236

5.15 4.30 5.21 6.04 10.8

1.38 0.93 0.95 3.10 2.13

4.27 4.11 4.92 8.17 5.21

0.810 0.850 0.940 1.70 1.09

2.38 2.37 2.70 5.03 2.95

0.363 0.368 0.391 0.748 0.449

2.31 2.34 2.52 4.88 2.85

0.330 0.369 0.405 0.786 0.448

5.01 5.62 7.31 2.82 4.89

1.68 1.67 2.31 1.14 1.59

5.64 6.11 7.71 10.8

0.940 1.11 1.29 1.78

2.40 2.70 3.12 4.05

0.323 0.357 0.382 0.502

2.13 2.18 2.35 3.03

0.312 0.306 0.327 0.416

8.00 8.36 8.76 7.88

5.87 5.32 6.51 6.86

1.18 0.840 0.630

5.18 3.76 2.94

1.43 1.11 0.860

0.416 0.398 0.241

1.62 1.32 0.980

0.310 0.240 0.180

1.97 1.50 1.24

0.460 0.340 0.280

1.34 1.04 0.900

0.221 0.165 0.164

1.49 1.16 1.24

0.231 0.190 0.223

15.7 11.8 14.5

2.05 1.60 2.22

9.06 7.24 9.48

2.45 1.93 2.11

0.782 0.608 0.547

2.56 2.11 2.26

0.460 0.360 0.430

2.90 2.24 2.89

0.630 0.500 0.670

1.94 1.44 2.12

0.299 0.243 0.353

2.04 1.61 2.49

0.337 0.259 0.409

55.6 37.2 129 29.2 86.8 138 203

5.87 3.57 15.3 3.49 9.98 15.8 22.8

8.73 3.40 11.4 2.83 7.28 11.9 16.3

1.71 0.620 1.72 0.490 1.25 1.96 2.71

10.3 3.78 9.54 3.12 7.43 11.2 16.2

1.88 0.710 1.87 0.680 1.56 2.20 3.19

5.04 1.92 5.17 2.03 4.70 6.55 9.45

0.728 0.276 0.759 0.322 0.731 0.998 1.41

4.71 1.72 4.85 2.18 4.86 6.81 9.23

0.731 0.258 0.803 0.357 0.803 1.10 1.44

22.2 12.1 62.2 13.9 37.7 61.0 86.7

5.80 2.46 12.9 2.85 8.24 12.9 17.8

1.08 1.04 2.39 1.40 1.14 2.00 1.90

9.59 15.1 10.0 10.2 10.5 10.7 11.4

Total Fe as Fe2 O3 , FeO* = Fe2 O3 * 0.9.

4. Analytical procedures 4.1. Major and trace elements Geochemical analyses for major elements were performed by XRF at the Center for Instrumental Analyses, Yamaguchi University. All analyses were made on glass beads using an alkali flux comprising lithium tetraborate. Trace elements including rare earth elements were determined with ICP-MS at Activation Laboratory Ltd., Canada using a lithium metaborate/tetraborate fusion technique. 4.2. U–Pb zircon age Uranium and Pb isotopes in zircon were analyzed using the sensitive high-resolution ion microprobe (SHRIMP II) at the National Institute of Polar Research, Tokyo (Japan). The analytical technique essentially follows Williams (1998). An O2 − primary ion beam of ∼2.1 and ∼5.1 nA was used to sputter an analytical spot of 20 and 25 ␮m diameter on the polished mount, respectively. TEMORA2 (206 Pb/238 U age: 416.78 ± 0.33 Ma; Black et al., 2004) and SL13 (U = 238 ppm) were used as standard materials for the U–Pb analysis. A correction for common Pb was made on the basis of the measured 204 Pb and the model for common Pb compositions proposed by Stacey and Kramers (1975). Data reduction and processing

were conducted using the computer programs SQUID 2 and ISOPLOT 3, provided by K.R. Ludwig at Berkeley Geochemistry Center, University of California (Ludwig, 2008, 2009). Analytical uncertainties listed in the data tables and plotted on the concordia diagrams are 1␴ and include measurement errors and uncertainties in the common Pb corrections. Uncertainties in the mean ages are 95% confidence limits. 4.3. Sr and Nd isotope geochemistry Isotopic analyses were performed with a thermal ionization mass spectrometer (MAT262) at Shimane University. The 87 Sr/86 Sr ratios and 143 Nd/144 Nd ratios were normalized to 86 Sr/86 Sr = 0.1194 and 146 Nd/144 Nd = 0.7219, respectively. The normalized 87 Sr/86 Sr ratios were corrected using the NBS-987 standard 87 Sr/86 Sr = 0.710241. The 143 Nd/144 Nd ratios were corrected with the Japanese standard JNdi-1 = 0.512106, which has been well characterized using the international standard La Jolla value of 143 Nd/144 Nd = 0.511849 (Tanaka et al., 2000). The Rb, Sr, Sm, and Nd concentrations were measured by inductively coupled plasma mass spectrometry (ICP-MS). The analytical errors for 87 Rb/86 Sr and 87 Sr/86 Sr were 1% (1␴) and 0.01% (1␴), respectively, and those for 147 Sm/144 Nd and 143 Nd/144 Nd ratios were 1% (1␴) and 0.01% (1␴), respectively. The initial Sr and Nd isotope ratios were calculated using the decay constants ␭87 Rb = 1.42 × 10−11 /y (Steiger and Jäger,

70

M. Owada et al. / Precambrian Research 234 (2013) 63–84

Fig. 4. Variation diagrams for the microgabbros and minette. (a and b) TAS and SiO2 versus K2 O diagrams. The compositional boundary between alkaline and non-alkaline rock series is quoted from Le Bas et al. (1986). (c) FeO*/MgO versus SiO2 diagram classified into tholeiite and calc-alkaline fields (Miyashiro, 1974). (d) FeO*/MgO versus TiO2 diagram. Note that the microgabbros are geochemically divided into low-Ti and high-Ti microgabbros. (e–h) FeO*/MgO versus Al2 O3 , CaO, Ni, Cr diagrams.

M. Owada et al. / Precambrian Research 234 (2013) 63–84

71

Fig. 5. Chondrite-normalized rare earth element patterns and spider diagrams for the low-Ti and high-Ti microgabbros (a and b), the migmatite and pelitic gneiss (c and d), the minette (e and f). Normalized values for chondrite and primitive mantle are quoted from Sun and McDonough (1989).

1977) and ␭147 Sm = 6.54 × 10−12 /y (Lugmair and Marti, 1978). The detailed isotopic analytical procedures were reported in Miyazaki and Shuto (1998). 5. Results 5.1. Major and trace elements The chemical compositions of the microgabbros, felsic to mafic schists, migmatites, and minettes are listed in Table 2 and Appendix 1. 5.1.1. Microgabbro, felsic schist, and mafic schist The microgabbro shows two modes of occurrence: (1) magmatic enclaves and (2) dikes. Both microgabbros have SiO2 contents ranging between 47 wt% and 53 wt% (Table 2) and belong to sub-alkaline, low-K, and tholeiite series signatures (Fig. 4a–c). The microgabbro of magmatic enclaves has low-Ti contents compared with that of dikes in the FeO*/MgO-versus-TiO2 diagram (Fig. 4d). We have therefore classified the microgabbros into low-Ti microgabbro and high-Ti microgabbro. Both microgabbros have Zr contents of less than 130 ppm (Table 2). The chemical

compositions of mafic schist resemble those of the low-Ti microgabbro (Fig. 4d). The microgabbros and mafic schist show constant values of Al2 O3 (Fig. 4e) and slight decrease in CaO (Fig. 4f) with FeO*/MgO ratios less than 2.0. On the other hand, Cr and Ni contents sharply decrease with increasing FeO*/MgO ratios (Fig. 4g and h). These chemical trends indicate olivine and probably chromite were the main fractionated phases, but clinopyroxene and plagioclase are not fractionated at the early stage of magmatic differentiation. The felsic schist shows more evolved compositions (Table 2). Zirconium contents of the mafic and felsic schists range from 20 to 40 ppm and from 40 to 120 ppm, respectively (Table 2). Chondrite-normalized rare earth element (REE) patterns of the microgabbros show slightly LREE enrichment with flat HREE (Fig. 5a). The trace element abundances of the low-Ti and high-Ti microgabbros, normalized to primitive mantle (Sun and McDonough, 1989), have almost flat patterns with negative Nb, Ta, and Ti anomalies (Fig. 5b) but both microgabbros show different patterns. The patterns of the low-Ti microgabbro show lower REE and trace element abundances than those of the high-Ti microgabbro (Fig. 5a and b). Moreover, the La/Nb ratios of the low-Ti microgabbro are higher than those of the high-Ti microgabbro (Table 2).

72

M. Owada et al. / Precambrian Research 234 (2013) 63–84

Fig. 6. Zircon CL and BSE images for dated samples. (a) Low-Ti microgabbro (CL, M00122602F), (b) High-Ti microgabbro (CL, M09011301B), (c) Felsic schist (CL, M09012401G), (d and e) Migmatite leucosome (CL, O90123101Ky2), (f) Minette (BSE, M09010802A). Circles marked on zircon grains indicate the measured points.

5.1.2. Kyanite-bearing migmatite Geochemically, the migmatite has a wide range of SiO2 contents, between 52 wt% (melanosome) and 74 wt% (leucosome). Pelitic gneisses associated with the migmatite were also measured and their SiO2 contents range from 52 wt% to 74 wt% (Table 2). Chondrite-normalized rare earth element patterns of the migmatite show enrichment of LREE and a depletion of HREE (Fig. 5c). The trace element abundances normalized to primitive mantle possess negative Nb, Ta, and Ti anomalies, indicating crustal signature (Fig. 5d). In addition, the patterns for migmatites show slightly convex shapes between Nd and Ti (Fig. 5d). 5.1.3. Minette The minettes range from 46 wt% to 54 wt% in SiO2 and are plotted within the field of alkaline rock and possesses high-K series signatures (Fig. 4a and b). The REE abundances of the minettes, normalized to the chondrite (Sun and McDonough, 1989), exhibit enrichment of LREE and depletion of HREE (Fig. 5e). A primitive mantle-normalized diagram (Sun and McDonough, 1989) is shown in Fig. 5f. The trace element abundances of the minettes display enrichment of LILE and depletion of HFSE with negative Nb, Ta, and Ti anomalies, similar to lamprophyres previously reported from the Sør Rondane Mountains (Ikeda et al., 1995). Some of the minettes have high MgO (>8 wt%) and Cr (>450 ppm) contents but relatively low Ni (<90 ppm) contents, indicating subtraction of olivine from a primary melt.

After polishing, analytical spots of samples and standard zircon were selected under backscattered electron (BSE) and cathodoluminescence (CL) imaging to assess their internal structure. Fig. 6 represents the selected BSE and CL images of zircon grains analyzed from samples. 5.2.1. Microgabbro (M08122602F, M09011301B) Only one zircon grain each were obtained from microgabbro samples (M08122602F; enclave, low-Ti microgabbro, M09011301B; dike, high-Ti microgabbro). These grains have euhedral shapes with 100 ␮m to 150 ␮m in diameter (Fig. 6a and b). The zircon grain from the microgabbro enclave contains mineral inclusions such as apatite and shows bright CL response domains at the edge (Fig. 6a), whereas the zircon of the microgabbro dike has oscillatory zoning. The apparent 206 Pb/238 U ages for the microgabbro enclave and dike are c. 998 Ma and c. 945 Ma, respectively (Table 3). These ages are correlated with magmatic ages of the tholeiitic tonalite (998–995 Ma) and the calc-alkaline tonalite (945–920 Ma), respectively, that coexist with the microgabbros (Table 1).

5.2. U–Pb zircon age

5.2.2. Felsic schist (M09012401G) The felsic schist contains prismatic zircons, ranging from 50 ␮m to 150 ␮m in diameter, with broad or irregular zoning surrounded by bright CL response domains (Fig. 6c). Thirty-three spots on 31 grains yield concordant U–Pb data, and the weighted mean of 206 Pb/238 U age is of 992 ± 4 Ma (MSWD = 0.92) (Table 4 and Fig. 7a). This age is interpreted as the magmatic age and is correlated with the older age-group of the tonalite (Table 1).

SHRIMP ages were obtained from zircon grains mounted in epoxy, which were separated from the microgabbro (M08122602F, M09011301B), felsic schist (M09012401G), kyanite-bearing leucosome of migmatite (O90123101Ky2), and minette (M09010802A).

5.2.3. Migmatite leucosome (O90123101Ky) Zircon grains in the leucosome occur in various shapes (rounded and well-rounded) and are up to 250 ␮m in grain size (Fig. 6d and e). The cores with oscillatory zoning show an age of 1238–633 Ma

1.46

5.2.4. Minette (M09010802A) The minette contains a few zircon grains with fragmented or prismatic shape and are up to 250 ␮m in size (Fig. 6f). These zircon grains are characterized by high U-contents (939–7053 ppm; Table 6). Fig. 7d presents a Terra–Wasserburg concordia plot of zircon age data (7 spots), having an age of 564 ± 2 Ma (MSWD = 0.92). This age is the same as the biotite K–Ar age of the minette within the analytical error (562 ± 14 Ma; Owada et al., 2010; Table 1). 5.3. Sr and Nd isotope geochemistry

Errors are 1-sigma; Pbc and Pb* indicate the common and radiogenic portions, respectively. Error in standard calibration was 0.19% (not included in above errors but required when comparing data from different mounts). Common Pb corrected using measured 204 Pb.

1.01 6.3 +1 920 ± 20 949 ± 30 945 ± 9 0.36 9 24 68 0.10

73

(206 Pb/238 U age: e.g., spots 41.2 and 42.1 in Table 5; Fig. 6e) with Th/U ratios of 0.04–0.30, whereas the rim of the grains with dark color in the CL-images range in age from 606 to 676 Ma (spots 22.1 and 17.1 in Table 5; Fig. 6d and e) with Th/U ratios of 0.17–0.21. No relationship was identified between 206 Pb/238 U age and Th/U ratios. Fig. 7b presents a Terra–Wasserburg concordia plot, displaying analyzed zircon age data (47 spots). The analytical results yield a concordia age of 620 ± 2 Ma (MSWD = 1.3). The analyzed data shows a wide range of inheritance between 700 and 1250 Ma (Fig. 7c). The concordia (620 Ma) and inherited (700–1250 Ma) ages are identical to the zircon U–Pb SHRIMP ages of the metamorphic rocks from the Sør Rondane Mountains reported in earlier studies (Table 1).

0.0707

0.38 0.0722 0.59 6.0 −1 993 ± 7 991 ± 8 0.51 763 0.02

M8122602F-1.1 low-Ti microgabbro M9011301B-1.1 high-Ti microgabbro

373

110

998 ± 5

% Dis cordant Pb/206 Pb Age

208 Pb/232 Th Age 207

Pb/238 U Age

206

Th/238 U 232

4-corr 206 Pb* (ppm) Th (ppm) U (ppm) Pbc (%) 206

Spot

Table 3 SHRIMP U–Pb zircon data from the low-Ti microgabbro (M08122602F) and the high-Ti microgabbro (M09011301B).

238

U/206 Pb*

±%

207

Pb*/ 206 Pb*

±%

M. Owada et al. / Precambrian Research 234 (2013) 63–84

Results of the Rb–Sr and Sm–Nd isotopic data are given in Table 7. The initial Nd and Sr ratios are calculated by the results of SHRIMP dating. The epsilon diagram is shown in Fig. 8. The initial epsilon Nd (␧NdI) values of the low-Ti and high-Ti microgabbros were +1 to +5, whereas the initial epsilon Sr (␧SrI) values of the low-Ti microgabbro are comparatively higher than those of the high-Ti microgabbro. The felsic and mafic schists are mostly plotted within the range of low-Ti microgabbro. The Sr–Nd isotopic compositions of the tonalites reported by Kamei et al. (2012) are also shown in this diagram (Fig. 8). The isotopic compositions of tonalites resemble those of the microgabbros and the mafic and felsic schists. Considering the petrography, geochemistry and geochronology, the mafic and felsic schists can be considered as highly deformed equivalents of the low-Ti microgabbro and the tholeiitic tonalite, respectively. The minette has limited isotopic compositions and is plotted close to the field of the microgabbros. The minettes give results of +7 to +10 for ␧SrI and −0.1 to +0.5 for ␧NdI. The migmatites and pelitic gneisses show higher radiogenic ␧SrI values and lower radiogenic ␧NdI values compared with the microgabbro, the felsic and mafic schists, and the minette (Table 7). 6. Discussion 6.1. Formation of Late Proterozoic microgabbros On the basis of the occurrence and geochronological data of the analyzed samples combined with previously reported age data (Table 1), the microgabbros and the tonalite represents coeval intrusive suites, therefore sharing the same tectonic setting. The low-Ti and high-Ti microgabbros belong to the non-alkaline and tholeiite series (Fig. 4a and c) and have almost flat patterns with the negative Nb, Ta, and Ti anomalies shown in the primitive mantle normalized spider diagram (Fig. 5b). These geochemical features suggest that the microgabbros were produced under a subductionrelated tectonic setting. Takahashi et al. (1990) suggested that in terms of its geochemical signatures, the tonalite was produced as a result of partial melting of an immature crust such as a primitive island arc. In addition, most of the zircon grains from the tonalite have euhedral and prismatic shapes, showing oscillatory zoning with no inherited

74

Table 4 SHRIMP U-Pb zircon data from the felsic schist (M09012401G). 206

M0912401G-1.1 M0912401G-1.2 M0912401G-2.1 M0912401G-3.1 M0912401G-4.1 M0912401G-5.1 M0912401G-6.1 M0912401G-7.1 M0912401G-8.1 M0912401G-9.1 M0912401G-10.1 M0912401G-11.1 M0912401G-12.1 M0912401G-13.1 M0912401G-13.2 M0912401G-14.1 M0912401G-15.1 M0912401G-16.1 M0912401G-17.1 M0912401G-18.1 M0912401G-19.1 M0912401G-20.1 M0912401G-21.1 M0912401G-22.1 M0912401G-23.1 M0912401G-24.1 M0912401G-25.1 M0912401G-26.1 M0912401G-27.1 M0912401G-28.1 M0912401G-29.1 M0912401G-30.1 M0912401G-31.1 M0912401G-32.1 M0912401G-33.1 M0912401G-34.1 M0912401G-35.1

– 0.00 – 0.14 0.05 0.26 0.03 0.08 0.14 0.10 0.22 – 0.08 – 0.02 – 0.04 0.00 0.33 0.14 – – 0.07 0.33 0.13 0.26 0.13 0.28 0.04 – 0.24 0.00 – – 0.05 0.15 0.15

Pbc (%)

U (ppm)

Th (ppm)

4-corr 206 Pb* (ppm)

232

Th/238 U

157 107 86 80 105 84 175 132 79 110 146 154 158 160 335 109 149 74 628 96 151 189 109 133 157 110 110 161 393 90 118 90 133 189 142 101 195

89 48 43 37 51 32 72 62 45 54 101 84 83 92 100 60 93 37 236 44 80 120 60 96 66 56 56 85 151 45 58 44 86 119 84 52 116

0.589 0.462 0.514 0.480 0.500 0.396 0.423 0.484 0.583 0.511 0.716 0.563 0.545 0.591 0.308 0.573 0.644 0.514 0.387 0.474 0.550 0.657 0.572 0.749 0.433 0.526 0.527 0.542 0.396 0.515 0.504 0.503 0.664 0.650 0.612 0.539 0.613

0.45 0.62 0.67 1.04 0.64 0.77 0.52 0.58 0.68 0.62 0.48 0.53 0.54 0.52 0.49 0.64 0.54 0.81 0.46 0.76 0.58 0.51 0.68 0.58 0.63 0.70 0.70 0.59 0.44 0.78 0.70 0.80 0.60 0.52 0.59 0.73 0.52

206

Pb/238 UAge

993 977 986 897 979 977 986 990 973 1008 994 982 1006 1000 1002 986 1012 990 938 1027 1006 958 981 1012 998 991 974 975 992 995 1006 996 1006 997 989 987 993

± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ±

9 11 12 17 12 13 10 11 13 12 10 10 11 10 9 12 11 15 9 14 11 10 13 13 11 13 13 11 8 14 13 15 12 11 12 14 11

207

Pb/206 PbAge

993 998 1033 943 981 960 974 999 988 978 968 1013 971 1017 1011 968 993 979 980 985 1021 989 995 916 949 925 960 895 1033 1048 915 1036 1017 1031 981 940 957

± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ±

Errors are 1-sigma; Pbc and Pb* indicate the common and radiogenic portions, respectively. Error in Standard calibration was 0.25% (not included in above errors but required when comparing data from different mounts). Common Pb corrected using measured 204 Pb.

19 23 31 52 27 41 19 24 35 28 30 22 24 23 15 100 23 30 29 36 23 22 30 42 29 41 35 37 15 33 40 28 23 21 26 38 27

208

Pb/232 ThAge

967 974 989 829 988 946 971 965 948 1005 948 1010 978 1022 1014 1005 1029 971 891 985 996 976 939 1017 957 967 954 901 1004 1011 958 961 995 975 978 967 965

± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ±

15 20 23 34 21 32 17 20 24 22 18 18 19 19 44 22 19 25 22 77 19 17 22 23 70 27 25 23 15 28 27 25 19 17 20 71 18

% Discordant −0 +2 +5 +5 +0 −2 −1 +1 +2 −3 −3 +3 −4 +2 +1 −2 −2 −1 +5 −5 +2 +3 +2 −11 −6 −8 −2 −10 +4 +5 −11 +4 +1 +4 −1 −5 −4

238

U/206 Pb*

6.00 6.11 6.05 6.70 6.10 6.11 6.05 6.03 6.14 5.91 6.00 6.08 5.92 5.96 5.95 6.05 5.88 6.02 6.39 5.79 5.92 6.24 6.09 5.88 5.97 6.02 6.13 6.12 6.01 5.99 5.92 5.99 5.92 5.98 6.03 6.05 6.00

±%

207

1.0 1.2 1.4 2.0 1.3 1.4 1.0 1.2 1.4 1.3 1.1 1.1 1.1 1.1 0.9 1.3 1.2 1.6 1.1 1.5 1.2 1.2 1.4 1.3 1.2 1.4 1.4 1.2 0.9 1.6 1.4 1.6 1.3 1.2 1.3 1.5 1.2

0.0722 0.0724 0.0737 0.0705 0.0718 0.0711 0.0716 0.0725 0.0721 0.0717 0.0714 0.0730 0.0715 0.0731 0.0729 0.0714 0.0722 0.0718 0.0718 0.0720 0.0733 0.0721 0.0723 0.0696 0.0707 0.0699 0.0711 0.0689 0.0737 0.0743 0.0696 0.0738 0.0731 0.0736 0.0718 0.0704 0.0710

Pb*/ 206 Pb*

±% 0.9 1.1 1.5 2.5 1.3 2.0 0.9 1.2 1.7 1.4 1.4 1.1 1.2 1.2 0.7 4.9 1.2 1.5 1.4 1.8 1.1 1.1 1.5 2.0 1.4 2.0 1.7 1.8 0.7 1.7 1.9 1.4 1.2 1.1 1.3 1.8 1.3

M. Owada et al. / Precambrian Research 234 (2013) 63–84

Spot

Table 5 SHRIMP U-Pb zircon data from the migmatite leucosome (O990123101Ky2). 206

O90123101Ky2-1.1 O90123101Ky2-1.2 O90123101Ky2-2.1 O90123101Ky2-2.2 O90123101Ky2-3.1 O90123101Ky2-4.1 O90123101Ky2-5.1 O90123101Ky2-5.2 O90123101Ky2-6.1 O90123101Ky2-7.1 O90123101Ky2-8.1 O90123101Ky2-9.1 O90123101Ky2-9.2 O90123101Ky2-10.1 O90123101Ky2-11.1 O90123101Ky2-11.2 O90123101Ky2-12.1 O90123101Ky2-13.1 O90123101Ky2-14.1 O90123101Ky2-14.2 O90123101Ky2-15.1 O90123101Ky2-16.1 O90123101Ky2-16.2 O90123101Ky2-17.1 O90123101Ky2-18.1 O90123101Ky2-18.2 O90123101Ky2-19.1 O90123101Ky2-20.1 O90123101Ky2-21.1 O90123101Ky2-22.1 O90123101Ky2-22.2 O90123101Ky2-25.1 O90123101Ky2-26.1 O90123101Ky2-27.1 O90123101Ky2-27.2 O90123101Ky2-28.1 O90123101Ky2-29.1 O90123101Ky2-30.1 O90123101Ky2-31.1 O90123101Ky2-23.1 O90123101Ky2-32.1 O90123101Ky2-33.1 O90123101Ky2-34.1 O90123101Ky2-35.1 O90123101Ky2-36.1 O90123101Ky2-36.2 O90123101Ky2-37.1 O90123101Ky2-37.2 O90123101Ky2-38.1 O90123101Ky2-39.1 O90123101Ky2-40.1 O90123101Ky2-41.1 O90123101Ky2-41.2

0.04 0.02 – 0.02 0.01 6.90 0.04 0.05 0.03 0.01 – 0.42 0.01 0.17 0.04 0.02 0.02 0.01 1.44 0.07 0.07 0.02 0.01 0.01 2.54 0.06 0.69 1.19 2.34 0.44 0.12 0.10 0.01 0.06 0.10 0.03 0.24 1.60 0.01 – 0.00 0.10 6.89 0.12 0.14 0.19 0.02 0.16 0.01 0.20 0.09 0.27 0.07

Pbc (%)

U (ppm)

Th (ppm)

4-corr 206 Pb* (ppm)

232

Th/238 U Age

206 Pb/238 U Age

159 468 2133 336 696 363 596 2148 2181 383 531 2044 364 2056 351 2126 1923 539 815 1904 86 215 567 578 217 2088 409 539 563 664 2296 2001 432 655 1966 547 328 480 2049 2439 2085 1913 540 230 496 2076 2329 1014 700 2401 439 789 241

56 81 83 124 118 125 128 78 111 117 117 108 139 100 135 72 101 131 96 60 22 76 122 99 140 104 122 124 105 136 122 60 130 117 71 114 99 79 81 179 116 78 102 48 176 42 106 288 133 110 147 81 72

24 40 189 29 60 31 50 194 192 34 45 162 53 178 52 186 169 45 72 166 7 26 49 51 34 190 35 47 48 56 202 172 37 58 172 48 38 62 183 210 190 166 66 32 83 180 208 159 106 215 57 70 44

0.36 0.18 0.04 0.38 0.18 0.36 0.22 0.04 0.05 0.32 0.23 0.05 0.40 0.05 0.40 0.04 0.05 0.25 0.12 0.03 0.26 0.36 0.22 0.18 0.67 0.05 0.31 0.24 0.19 0.21 0.05 0.03 0.31 0.18 0.04 0.22 0.31 0.17 0.04 0.08 0.06 0.04 0.20 0.21 0.37 0.02 0.05 0.29 0.20 0.05 0.35 0.11 0.31

1037 607 633 608 617 606 597 646 628 628 608 568 1008 619 1018 624 627 598 628 624 610 851 612 625 1075 649 616 620 615 606 628 614 610 628 624 623 817 907 638 615 649 621 864 957 1153 621 637 1080 1045 639 906 634 1235

± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ±

34 12 5 14 5 5 9 5 5 16 11 4 32 5 9 5 9 11 5 8 7 9 10 5 10 19 5 5 11 11 9 5 10 5 5 5 25 61 10 6 32 5 56 37 26 5 9 17 34 10 8 21 11

207

Pb/206 Pb Age

1031 604 629 630 613 660 630 641 611 627 618 625 1055 616 1063 619 610 603 689 616 612 921 603 676 1108 642 622 632 772 617 637 618 578 626 599 618 849 914 608 609 614 623 941 965 1113 607 623 1078 1072 642 988 617 1250

± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ±

87 17 8 21 14 102 16 8 8 20 16 67 14 11 15 9 9 18 85 10 48 21 17 76 78 37 40 71 120 25 53 39 19 17 11 18 22 204 8 32 8 52 220 186 41 12 29 10 68 10 15 20 110

208

Pb/232 Th Age

1005 598 622 596 568 624 580 646 625 628 612 594 1028 604 1026 590 603 573 636 587 609 871 608 585 1084 631 692 641 709 616 639 629 598 638 624 609 1398 870 608 616 648 623 862 969 1181 756 626 1071 1034 645 1014 651 1159

± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ±

85 14 9 29 8 44 11 27 10 29 13 31 50 21 13 13 13 12 64 22 20 14 12 9 40 23 18 22 67 29 18 26 12 34 24 9 70 76 13 9 42 22 105 44 30 92 26 20 60 25 28 40 19

% Discordant −1 −1 −1 +4 −1 +9 +5 −1 −3 −0 +2 +10 +5 −0 +5 −1 −3 +1 +9 −1 +0 +8 −2 +8 +3 −1 +1 +2 +21 +2 +1 +1 −6 −0 −4 −1 +4 +1 −5 −1 −6 +0 +9 +1 −4 −2 −2 −0 +3 +0 +9 −3 +1

238

U/206 Pb*

5.73 10.12 9.69 10.11 9.96 10.15 10.30 9.49 9.77 9.78 10.10 10.86 5.91 9.92 5.84 9.84 9.79 10.28 9.77 9.83 10.07 7.09 10.04 9.82 5.51 9.44 9.97 9.90 9.98 10.15 9.78 10.00 10.08 9.77 9.84 9.85 7.41 6.62 9.61 9.99 9.44 9.89 6.98 6.25 5.10 9.89 9.63 5.49 5.68 9.60 6.63 9.68 4.74

±%

207

3.6 2.1 0.8 2.5 0.8 0.9 1.6 0.8 0.8 2.6 1.9 0.8 3.4 0.8 0.9 0.8 1.6 1.8 0.8 1.3 1.3 1.1 1.7 0.8 1.0 3.0 0.9 0.9 1.9 1.8 1.4 0.8 1.8 0.8 0.8 0.9 3.2 7.3 1.6 1.1 5.2 0.8 7.0 4.2 2.5 0.8 1.5 1.7 3.5 1.6 0.9 3.5 1.0

0.0736 0.0600 0.0607 0.0607 0.0603 0.0616 0.0607 0.0610 0.0602 0.0607 0.0604 0.0606 0.0745 0.0604 0.0748 0.0604 0.0602 0.0600 0.0624 0.0603 0.0602 0.0698 0.0600 0.0621 0.0765 0.0611 0.0605 0.0608 0.0649 0.0604 0.0609 0.0604 0.0593 0.0606 0.0599 0.0604 0.0674 0.0695 0.0601 0.0601 0.0603 0.0605 0.0704 0.0713 0.0767 0.0601 0.0605 0.0754 0.0751 0.0611 0.0721 0.0604 0.0822

Pb*/ 206 Pb*

±% 4.3 0.8 0.4 1.0 0.6 4.7 0.8 0.4 0.4 0.9 0.8 3.1 0.7 0.5 0.7 0.4 0.4 0.8 4.0 0.5 2.2 1.0 0.8 3.5 3.9 1.7 1.8 3.3 5.7 1.2 2.5 1.8 0.9 0.8 0.5 0.8 1.1 9.9 0.4 1.5 0.4 2.4 10.8 9.1 2.1 0.5 1.3 0.5 3.4 0.5 0.7 0.9 5.6

M. Owada et al. / Precambrian Research 234 (2013) 63–84

Spot

75

76

Table 5 (Continued) Spot

Pbc (%)

– 4.08 0.01 1.95 – – 0.08 0.00 0.04 0.03 0.01 2.11 1.47 0.08 0.03 0.07 16.85 7.54 0.81 0.07 0.02 0.51 0.03 0.07 7.23 0.04 1.18 0.04 1.63 5.32 18.06 3.44 6.18 – 0.35 0.02 0.20 0.05 1.52

U (ppm)

Th (ppm)

4-corr 206 Pb* (ppm)

232

Th/238 U Age

206

Pb/238 U Age

358 481 2093 211 823 356 430 1685 2164 482 555 629 500 381 321 710 101 525 591 344 316 231 785 569 711 522 359 532 534 714 540 582 202 323 384 801 93 534 474

131 106 200 74 110 98 110 158 115 86 102 222 123 118 138 126 30 122 113 104 98 41 128 108 139 125 163 123 60 113 109 109 64 119 163 118 38 119 107

51 40 186 24 76 31 38 186 191 52 48 71 43 29 51 61 19 45 54 30 28 23 77 49 87 44 58 46 65 61 47 49 28 28 58 68 11 46 45

0.38 0.23 0.10 0.36 0.14 0.28 0.26 0.10 0.05 0.18 0.19 0.36 0.25 0.32 0.44 0.18 0.31 0.24 0.20 0.31 0.32 0.18 0.17 0.20 0.20 0.25 0.47 0.24 0.12 0.16 0.21 0.19 0.33 0.38 0.44 0.15 0.43 0.23 0.23

987 596 634 805 658 622 629 779 631 769 619 799 615 555 1097 618 1298 610 652 620 644 721 700 613 860 608 1109 622 856 610 626 605 950 612 1040 609 825 615 674

± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ±

35 5 15 22 9 5 17 18 8 6 5 7 14 5 55 5 88 13 12 15 6 7 14 11 14 16 11 18 7 5 11 5 25 5 9 5 10 5 6

207 Pb/206 Pb Age

1090 613 622 923 687 595 643 839 612 768 596 901 674 593 1023 607 1544 800 741 607 638 690 721 617 1017 611 1063 593 912 732 569 679 1098 611 1003 606 769 596 652

Errors are 1-sigma; Pbc and Pb* indicate the common and radiogenic portions, respectively. Error in Standard calibration was 0.19% (not included in above errors but required when comparing data from different mounts). Common Pb corrected using measured 204 Pb.

± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ±

89 79 8 125 12 21 21 57 9 16 17 142 45 26 99 16 129 388 83 107 22 38 13 19 160 19 91 19 34 161 412 62 133 24 20 15 44 19 76

208 Pb/232 Th Age

982 621 640 715 702 640 625 986 627 719 605 922 639 536 1137 584 1809 705 704 604 636 726 690 605 974 610 1132 616 836 718 504 664 1093 617 1015 604 826 605 636

± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ±

49 49 16 46 12 10 19 40 13 11 9 74 35 9 61 9 271 84 26 17 10 34 16 15 86 18 67 20 65 78 108 52 85 9 15 9 22 25 50

% Discordant +10 +3 −2 +14 +4 −5 +2 +8 −3 −0 −4 +12 +9 +7 −8 −2 +18 +25 +13 −2 −1 −5 +3 +1 +17 +0 −5 −5 +7 +17 −11 +11 +14 −0 −4 −0 −8 −3 −4

238

U/206 Pb*

6.05 10.32 9.68 7.52 9.30 9.86 9.75 7.78 9.73 7.89 9.92 7.58 10.00 11.12 5.39 9.94 4.48 10.07 9.39 9.90 9.52 8.46 8.72 10.03 7.01 10.11 5.33 9.87 7.04 10.08 9.80 10.16 6.30 10.05 5.71 10.10 7.32 9.99 9.07

±%

207

3.8 0.9 2.5 2.9 1.4 0.9 2.8 2.5 1.4 0.9 0.9 1.0 2.3 0.9 5.5 0.8 7.5 2.3 2.0 2.5 0.9 1.0 2.0 2.0 1.7 2.7 1.0 3.0 0.9 0.9 1.9 0.9 2.8 0.9 0.9 0.8 1.2 0.9 1.0

0.0758 0.0603 0.0605 0.0698 0.0624 0.0598 0.0611 0.0670 0.0602 0.0648 0.0598 0.0691 0.0620 0.0597 0.0733 0.0601 0.0958 0.0658 0.0640 0.0601 0.0610 0.0625 0.0634 0.0604 0.0731 0.0602 0.0748 0.0597 0.0694 0.0637 0.0590 0.0622 0.0761 0.0602 0.0726 0.0601 0.0648 0.0598 0.0614

Pb*/ 206 Pb*

±% 4.4 3.7 0.4 6.1 0.6 1.0 1.0 2.7 0.4 0.7 0.8 6.9 2.1 1.2 4.9 0.8 6.8 18.5 3.9 5.0 1.0 1.8 0.6 0.9 7.9 0.9 4.5 0.9 1.7 7.6 18.9 2.9 6.7 1.1 1.0 0.7 2.1 0.9 3.5

M. Owada et al. / Precambrian Research 234 (2013) 63–84

O90123101Ky2-42.1 O90123101Ky2-42.2 O90123101Ky2-43.1 O90123101Ky2-44.1 O90123101Ky2-45.1 O90123101Ky2-46.1 O90123101Ky2-48.1 O90123101Ky2-49.1 O90123101Ky2-50.1 O90123101Ky2-51.1 O90123101Ky2-52.1 O90123101Ky2-53.1 O90123101Ky2-54.1 O90123101Ky2-55.1 O90123101Ky2-56.1 O90123101Ky2-57.1 O90123101Ky2-58.1 O90123101Ky2-59.1 O90123101Ky2-60.1 O90123101Ky2-61.1 O90123101Ky2-62.1 O901231011Ky2-63.1 O90123101Ky2-64.1 O90123101Ky2-65.1 O90123101Ky2-66.1 O90123101Ky2-67.1 O90123101Ky2-68.1 O90123101Ky2-69.1 O90123101Ky2-70.1 O90123101Ky2-71.1 O90123101Ky2-72.1 O90123101Ky2-73.1 O90123101Ky2-75.1 O90123101Ky2-76.1 O90123101Ky2-77.1 O90123101Ky2-78.1 O90123101Ky2-79.1 O90123101Ky2-80.1 O90123101Ky2-81.1

206

0.53 3.23 11 5.90 2.87 0.41 0.82 0.35 0.29 0.31 0.18 0.39 0.0595 0.0589 0.0596 0.0594 0.0582 0.0585 0.0588 0.0588 0.0587 0.0590 0.0586 0.0586 0.55 0.56 0.78 0.67 0.58 0.55 0.54 0.54 0.55 0.54 0.54 1.18 10.9 11.2 9.9 10.9 11.0 10.6 9.9 10.9 10.9 11.0 9.5 10.9 +3 +2 −5 +2 −4 −7 −12 −1 −1 +1 −17 −3 6 14 461 66 24 4 33 6 3 4 5 7 ± ± ± ± ± ± ± ± ± ± ± ± Errors are 1-sigma; Pbc and Pb* indicate the common and radiogenic portions, respectively. Error in Standard calibration was 0.19% (not included in above errors but required when comparing data from different mounts). Common Pb corrected using measured 204 Pb.

581 559 695 565 551 570 601 555 551 554 623 536 12 70 237 128 63 9 18 8 6 7 4 9 ± ± ± ± ± ± ± ± ± ± ± ± 585 562 589 581 539 548 560 560 557 566 554 551 3 3 5 4 3 3 3 3 3 3 3 6 ± ± ± ± ± ± ± ± ± ± ± ± 566 553 620 568 559 583 622 565 565 562 643 566 0.23 0.33 0.05 0.26 0.33 0.38 0.05 1.04 0.32 0.37 0.04 0.30 210 162 611 273 73 181 464 250 227 314 537 195 589 668 359 857 299 813 275 3196 906 1447 225 709 0.52 6.55 3.20 1.94 1.27 0.21 0.08 0.18 0.01 0.18 0.00 0.15 M9010802A-1.1 M9010802A-2.1 M9010802A-3.1 M9010802A-4.1 M9010802A-5.1 M9010802A-6.1 M9010802A-7.1 M9010802A-8.1 M9010802A-9.1 M9010802A-10.1 M9010802A-11.1 M9010802A-12.1

2663 2106 7053 3445 939 2229 5332 3175 2880 4015 5950 2477

Th (ppm) U (ppm) Pbc (%) 206

Spot

Table 6 SHRIMP U–Pb zircon data from the minette (M09010802A).

4-corr 206 Pb* (ppm)

232

Th/238 U

206

Pb/238 U Age

207 Pb/206 Pb Age

208

Pb/232 Th Age

% Discordant

238

U/206 Pb*

±%

207

Pb*/ 206 Pb*

±%

M. Owada et al. / Precambrian Research 234 (2013) 63–84

77

cores (Shiraishi et al., 2008; Kamei et al., 2012). On the basis of the geochemical and geochronological features, the tonalite complex would have been produced within a volcanic arc, probably in a juvenile arc setting. The low-Ti microgabbro is geochemically characterized by a wide range of ␧SrI values (Fig. 8). Field observations suggest that the low-Ti microgabbro has interacted with the host tonalite magma. Sample M08122602F, however, represents geochemical features very close to primary compositions because FeO*/MgO ratio, Cr content, and Ni content are c. 1.0, 560 ppm, and 140 ppm, respectively, although the ␧SrI value of this sample is +8.01 (Tables 2 and 7). Other samples, M08122602A and M09011101, are also similar to the primary basaltic compositions, with ␧SrI values of +9.21 and −14.3, respectively. These ␧SrI values cover the compositional range of the low-Ti microgabbro. Furthermore, most ␧SrI values of the host tonalite are in the range between +15 and −25 (Fig. 8; Takahashi et al., 1990; Owada et al., 2006; Kamei et al., 2012), similar to those of the low-Ti microgabbro. Therefore, the ␧SrI values of the low-Ti microgabbro strongly reflect enriched Sr isotopic compositions of a source mantle, although isotopic modification due to interaction with the host tonalite cannot be ruled out. Sr–Nd isotopic systematics and trace element compositions reveal that the low-Ti and high-Ti microgabbros are derived from different mantle sources. The low-Ti microgabbro defines a broad negative trend in ␧SrI and ␧NdI data (Fig. 8). These data may be most easily accounted for by involving subduction components with high ␧SrI and low ␧NdI values in the petrogenesis of the magmas. On the other hand, the high-Ti microgabbro shows the same range of ␧NdI values but limited ␧SrI values compared to the low-Ti microgabbro (Fig. 8). Fig. 9 shows Ba/Nb ratios versus ␧NdI values for the low-Ti and high-Ti microgabbros. The Ba/Nb ratios of the low-Ti microgabbro are higher than those of the high-Ti microgabbro and the ␧NdI values of both microgabbros are overlapped. Since Ba can be easily mobilized by a fluid phase rather than HFSE including Nb and REEs (Tatsumi et al., 1984, 1986), the low-Ti microgabbro tends to be influenced by hydrous fluids, probably derived from the subducted oceanic slab. In the V (ppm) versus Ti/100 diagram (Fig. 10), the low-Ti and high-Ti microgabbros are plotted within the fields of the arc tholeiite and MORB and back-arc basin basalts, respectively. Compared to N-type MORB, the back-arc basalts show relative enrichment in the LILE, the mobile elements, which are transported into the source of island-arc basalts by subduction-zone fluids (Wilson, 1989). The high-Ti microgabbro shows slight evidence of negative Nb anomalies with LILE enrichment compared to MORB (Fig. 5b). These geochemical features seem to be matched with the back-arc basalt rather than MORB. Most of the low-Ti microgabbros belong to the oceanic arc field in Zr/Y versus Zr ppm diagram (Fig. 11a). On the other hand, the high-Ti microgabbros are plotted within the range of the back-arc basin basalts in oceanic arc settings, but not in continental arc settings (Fig. 11b). A plausible tectonic model relating to this event is shown in Fig. 12. The SHRIMP results date the low-Ti microgabbro as being older than the high-Ti microgabbro (Table 3). Considering the geochemical features and geochronological results, the early Neoproterozoic tonalite complex in the Sør Rondane Mountains would be formed firstly by subduction-related magmatism as a juvenile arc setting at c. 998 Ma, and subsequently intruded by back-arc basaltic magma at c. 948 Ma. Such magmatism probably progressed in an oceanic arc setting. 6.2. Timing of metamorphism and its tectonic background The zircon U–Pb SHRIMP dating for the migmatite leucosome gives an age of 620 Ma (Fig. 7b and c). This age can be correlated to the main metamorphic event proposed by Shiraishi et al.

78

M. Owada et al. / Precambrian Research 234 (2013) 63–84

Fig. 7. Terra – Waserburg concordia diagrams. (a) Felsic gneiss, (b) Minette, (c and d) Migmatite leucosome.

Fig. 9. Ba/Nb versus ␧NdI diagram for the low-Ti and high-Ti microgabbros.

Fig. 8. Epsilon diagram for the low-Ti and high-Ti microgabbros, the mafic and felsic schists, and the minette. Sr and Nd isotopic ratios are corrected with magmatic ages for each rock type. Compositional fields of the coexisting tonalite (tholeiitic and calc-alkaline) are also shown. The tonalites data are taken from Kamei et al. (2012).

(2008). In a Gondwanan context, Shiraishi et al. (2008) summarized the tectonothermal histories of the Sør Rondane Mountains as follows. The Sør Rondane Mountains were metamorphosed under granulite-facies conditions at 640–600 Ma, then retrograde amphibolite-facies conditions at c. 570 Ma. The high-grade metamorphism was followed by the extensive A-type granitoid activity and contact metamorphism, between 560 and 500 Ma. The timing

M. Owada et al. / Precambrian Research 234 (2013) 63–84

79

Fig. 12. Tectonic model for the formation of high-Ti microgabbro at c. 950 Ma. The high-Ti microgabbro was produced at a back-arc setting and intruded the oceanicarc crust comprising the tonalite and the low-Ti microgabbro.

Fig. 10. Ti/100 versus V diagram (Shervais, 1982) for the low-Ti and high-Ti microgabbros.

of peak metamorphism during 640–600 Ma corresponds with the early stage of the Pan-African suture event, and the following igneous activity (560–500 Ma) reflects the extensional collapse of the Orogen. The highest-grade gneisses are located in the central- to northeastern parts of the Sør Rondane Mountains, and their peak metamorphic temperatures are up to 900 ◦ C (Asami et al., 2007; Adachi et al., 2010; Nakano et al., 2011; Baba et al., 2012). The migmatite sample dated here was collected from the highest-grade metamorphic region, Brattnipene, in the central part of the Sør Rondane Mountains (Fig. 1). Asami and Shiraishi (1987), Asami et al. (1992) described retrograde kyanite in the pelitic gneiss from the Koyubi-ridge of Brattnipane (Fig. 1). The sample was collected from nearby locality to the retrograde kyanite-bearing pelitic gneiss as described by Asami and Shiraishi (1987) and therefore, the formation of the migmatite leucosome is considered to have occurred

during the retrograde metamorphism, thereby dating the peak metamorphism prior to 620 Ma, most probably between 640 and 620 Ma. The Sør Rondane Mountains has been divided into NE-terrane and SW-terrane by a fault and/or thrust in terms of their progressive metamorphism and P–T paths (Osanai et al., 2013). Since the timing of peak metamorphism was caused by the initial stage of the Pan-African suture event (Shiraishi et al., 2008), this metamorphic event prior to 620 Ma is assumed to have occurred during the amalgamation between NE-terrane and SW-terrane, related to the formation of the Gondwana supercontinent (Osanai et al., 2013; Baba et al., 2012). 6.3. Geochemical characteristics of the minettes The minettes described here are important for understanding the chemical characteristics of the source mantle, since some of the minettes show high MgO and Cr contents and low FeO*/MgO ratios, indicating magmas that are slightly evolved but remain close to primary compositions. The minette magma shows high K, Rb, and Ba contents. In general, a mantle-derived magma produced by low degree of melting has a high concentration of incompatible elements such as LILE. The minettes with MgO contents greater

Fig. 11. Discrimination Zr versus Zr/Y diagram (Pearce, 1983) for the low-Ti microgabbro (a). Note: The low-Ti microgabbro is plotted within the oceanic arc field. Nb/Yb versus Th/Yb diagram for the high-Ti microgabbro (b). Note: The diagram plots the compositions of back-arc basin basalts situated in continental arcs and oceanic arcs. The high-Ti microgabbro belongs to the range of oceanic arc settings. Compiled data for back-arc basin basalts are taken from the East Scotia ridge (oceanic arc; Fretzdorff et al., 2002), the Maus, Papua New Guinea (oceanic arc; Sinton et al., 2003) and the Loncopue graben, Argentina (continental arc; Varekamp et al., 2010).

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Table 7 Sr and Nd isotopic composions of the microgabbros, minette and metamorphic rocks in the Sør Rondane Mountains. Sample

Sr (ppm)

87

Rb/86 Sr

87

Sr/86 Sr

2␴

SrI

␧Sr

Nd (ppm)

Sm (ppm)

147

Sm/144 Nd

143

Nd/144 Nd

2␴

NdI

␧NdI

Age (Ma)

5 6 1 2 33

146 134 134 181 428

0.0991 0.1295 0.0216 0.0320 0.2230

0.70549 0.70446 0.70425 0.70432 0.70545

1 1 1 1 1

0.70409 0.70263 0.70395 0.70387 0.70229

11.2 −9.56 9.21 8.08 −14.3

6.11 8.69 3.85 16.5 16.6

1.65 3.07 1.17 3.81 3.07

0.1633 0.2136 0.1837 0.1396 0.1118

0.512542 0.512931 0.512697 0.512364 0.512331

13 12 14 14 14

0.511482 0.511544 0.511504 0.511457 0.511605

2.36 3.57 2.79 1.88 4.77

990 990 990 990 990

Microgabbro (high-Ti) M08122101F 7 M08122601B 40 114 M08122602D 5 M09010702 12 M09011301B

299 172 378 108 285

0.0677 0.6731 0.8732 0.1339 0.1218

0.70368 0.71157 0.71533 0.70488 0.70431

1 1 1 1 1

0.70276 0.70242 0.70347 0.70306 0.70266

−9.06 −13.9 1.03 −4.84 −10.5

13.9 14.6 24.9 14.7 16.0

3.89 3.90 5.83 5.57 4.52

0.1692 0.1615 0.1415 0.2291 0.1708

0.512704 0.512669 0.512522 0.512902 0.512746

12 14 14 14 14

0.511650 0.511663 0.511640 0.511474 0.511682

4.63 4.88 4.43 1.19 5.25

950 950 950 950 950

147 236 224 264

660 631 2030 1790

0.6445 1.0827 0.3192 0.4267

0.70954 0.71306 0.70700 0.70774

1 1 1 1

0.70440 0.70442 0.70445 0.70434

8.50 8.77 9.23 7.60

56.6 53.7 73.4 94.8

0.1335 0.1306 0.1260 0.1326

0.512416 0.512391 0.512407 0.512416

12 12 13 14

0.511926 0.511912 0.511945 0.511929

0.18 −0.10 0.54 0.24

560 560 560 560

Mafic schist (inclusion) M08122602B M09012401C M08012401J

12 2 4

143 195 178

0.2428 0.0297 0.0650

0.70860 0.70437 0.70501

1 1 1

0.70517 0.70395 0.70408

26.5 9.21 11.2

5.18 3.76 2.94

1.43 1.11 0.86

0.1669 0.1785 0.1768

0.512511 0.512638 0.512640

13 13 13

0.511427 0.511479 0.511492

1.29 2.30 2.55

990 990 990

Felsic schist (inclusion) M09012401G M09012401K1 M09012401L

3 1 12

179 161 138

0.0485 0.0180 0.2516

0.70498 0.70342 0.70627

1 1 1

0.70429 0.70316 0.70271

14.1 −1.92 −8.41

9.06 7.24 9.48

2.45 1.93 2.11

0.1635 0.1612 0.1346

0.512520 0.512570 0.512474

14 14 14

0.511458 0.511523 0.511600

1.90 3.17 4.67

990 990 990

Migmatite & pelitic gneiss 180 O90123101Ky1-2 122 O90122101Ky2 63 M08122801E 37 M09011602F M09012301H 119

233 290 288 84 61

2.2391 1.2184 0.6332 1.2762 5.6872

0.72579 0.71772 0.71297 0.72210 0.78547

1 1 1 1 1

0.70598 0.70695 0.70737 0.71082 0.73518

5.80 2.46 12.9 2.85 8.24

0.1579 0.1229 0.1254 0.1239 0.1321

0.512362 0.512283 0.512353 0.512409 0.512146

13 12 14 11 14

0.511720 0.511784 0.511844 0.511906 0.511609

−2.32 −1.08 0.10 1.30 −4.50

620 620 620 620 620

Minette M08121201B M08121201C M09010801 M09010802A

31.0 44.7 50.7 99.7 446

22.2 12.1 62.2 13.9 37.7

12.5 11.6 15.3 20.8

M. Owada et al. / Precambrian Research 234 (2013) 63–84

Microgabbro (low-Ti) M09010703A M08122301D M08122602A M08122602F M09011101

Rb (ppm)

3.5 14 15 11 12

Nd (ppm)

5.20 6.04 10.78 5.01 5.61 Trace element compositions of each primary melt are calculated by a fomulation that utilizes Rayleigh fractionation model with partitioning coefficient of olivine. Partitioning coefficient of olivine is quoted from Fujimaki et al. (1984), McKenzie and O’Nions (1991) and Brunet and Chazot (2001).

Ce/Sm (ppm) Rb/Sr (ppm)

0.0075 0.011 0.093 0.023 0.286 0.18 0.18 0.21 0.26 0.31

Lu (ppm) Yb (ppm)

1.1 1.1 1.3 1.8 1.9 0.17 0.17 0.19 0.29 0.30

Tm (ppm) Er (ppm)

1.0 1.1 1.3 1.9 2.0 0.35 0.41 0.45 0.64 0.70

Ho (ppm)

1.6 2.0 2.2 3.4 3.4

Dy (ppm)

0.24 0.37 0.37 0.57 0.58

Tb (ppm) Gd (ppm)

1.3 2.5 2.4 3.1 3.4 0.38 0.62 1.2 1.0 1.1

Eu (ppm)

Low-Ti Low-Ti Low-Ti High-Ti High-Ti M08122602A M08122602F M09011101 M08122101F M08122601B

Sm (ppm) Rock type Sample

1.1 3.3 2.7 3.1 3.2

0.75 3.0 3.8 2.3 2.5 5.6 20 29 15 18 2.6 7.2 11 6.7 7.5 9.3 11 12 19 19 0.46 1.8 0.89 1.8 1.8 17 67 27 74 71 123 157 344 236 123 0.018 0.15 0.071 0.24 0.30 0.18 1.8 1.6 4.6 4.0 Low-Ti Low-Ti Low-Ti High-Ti High-Ti M08122602A M08122602F M09011101 M08122101F M08122601B

0.92 1.7 32 5.5 35

29 59 163 57 228

0.75 4.0 0.51 0.57 0.91

Sr (ppm) Rock type Sample

Rb (ppm)

Ba (ppm)

Th (ppm)

Nb (ppm)

Ta (ppm)

Zr (ppm)

Hf (ppm)

Y (ppm)

La (ppm)

Ce (ppm)

Pr (ppm)

0.69 0.70 0.70 0.69 0.70 100.00 100.00 100.00 100.00 100.00

Total (wt%) P2 O5 (wt%)

0.05 0.04 0.04 0.08 0.10 0.09 0.21 1.11 0.31 0.94

K2 O (wt%) Na2 O (wt%)

3.10 2.14 2.78 1.83 2.14 11.44 10.98 9.15 8.29 7.99

CaO (wt%)

0.17 0.25 0.20 0.15 0.17 7.01 9.81 9.50 10.88 10.91 14.40 13.13 13.46 14.30 13.69

MgO (wt%) MnO (wt%) FeO (wt%) Al2 O3 (wt%)

0.38 0.48 0.58 1.14 1.11

TiO2 (wt%) SiO2 (wt%)

53.23 48.92 49.51 47.34 47.32 8.0 13.5 11.0 21.5 22.0 Low-Ti Low-Ti Low-Ti High-Ti High-Ti M08122602A M08122602F M09011101 M08122101F M08122601B

Added olivine % Rock type Sample

Table 8 Estimated primary basaltic melts of the low-Ti and high-Ti microgabbros.

10.13 14.02 13.65 15.67 15.63

FeO/MgO (wt%)

M. Owada et al. / Precambrian Research 234 (2013) 63–84

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than 8 wt% show high Ce/Sm and Sm/Yb ratios (Table 2), indicating that the melting rates are low (Hawksworth et al., 1994). The Ba contents of minettes are up to 9000 ppm with 8 wt% of MgO (Table 2). The primitive mantle has Ba content of c. 7 ppm (Sun and McDonough, 1989). To obtain the Ba content of the minette magma derived from primitive mantle compositions, the inferred degree of melting is 0.3–0.08% if the bulk partitioning-coefficient of Ba is close to zero. These low degrees of melting are unlikely to cause segregation of the mafic magmas from the source mantle. Therefore, the minette magmas would be originated from metasomatized mantle with initial enrichment in LILE. The minettes show negative Nb, Ta, and Ti anomalies and the LILE/HFSE ratios are higher than those of Oceanic Island Basalts (OIB) (Fig. 5c). Such geochemical features resemble the previously reported lamprophyres and high-K mafic intrusive rocks from central to eastern Dronning Maud Land (Arima and Shiraishi, 1993; Ikeda et al., 1995; Zhao et al., 1995; Hoch et al., 2001; Mikhalsky, 2004; Owada et al., 2008, 2010). These lamprophyres and high-K mafic intrusive rocks are plotted in the within-plate field and a part of the island arc field that is close to the within-plate field on some discrimination diagrams (Ikeda et al., 1995; Owada et al., 2008). On the basis of these geochemical features, Ikeda et al. (1995) interpreted that the lamprophyre from the Sør Rondane Mountains was formed in a within-plate tectonic setting by the mixing of subduction-related materials (e.g., slab-derived fluids, melting product of subducted crustal rocks, or reaction with fossil wedge mantle) at mantle depth. The ␧SrI and ␧NdI values of the minettes show a limited range between the microgabbros and the pelitic gneiss and migmatite (Table 7). Moreover, the timing of magma activity of the minette occurs several tens million years later than the peak metamorphism, at which NE-terrane and SW-terrane of the Sør Rondane Mountains would have amalgamated. The petrogenesis of minette magma should, therefore, be thought of within the context of lithospheric evolution beneath the Sør Rondane Mountains, from the end of the Mesoproterozoic (1000 Ma) until the episodic collision event (650–620 Ma). 6.4. Evolution of lithospheric mantle underneath the Sør Rondane Mountains To assess the lithospheric evolution, it is important to determine the primary basalt compositions during the early Neoproterozoic (c. 1000 Ma) because such melt should geochemically reflect the source mantle signature. The low-Ti and high-Ti microgabbros with primary compositions are rare, but some samples are useful to estimate the primary melt compositions coexisting with a source mantle. Both microgabbros show chemical trends of increasing FeO*/MgO ratios with immediately decreasing Ni contents, indicating olivine fractionation (Fig. 4f). Therefore, it is possible to estimate the primary melt composition using the olivine maximum fractionation method of Tatsumi et al. (1983). Five samples are candidates for the estimation of primary melt compositions, three low-Ti microgabbros (M082602A, M08122602F, M09011101) and two high-Ti microgabbros (M08122101F, M08122601B). The procedure for calculating the primary basaltic compositions is adapted from Tatsumi et al. (1983) and Tamura et al. (2011). Assumptions in this model are as follows: (1) Fe–Mg exchange partitioning between olivine and liquid, KD = (Fe/MgOl /Fe/Mgliq ) = 0.3, (2) (Fe2+ /Fe2+ +Fe3+ ) magma = 0.9, (3) mantle olivine Fo = 89.5. The chemical compositions of magmas are estimated by the following methods. The olivine composition in equilibrium with the basaltic magma is estimated using Fe–Mg exchange partitioning. Then, 0.5 wt% of the equilibrium olivine is added to the magma. This process is repeated until the magma finally equilibrates with Fo = 89.5

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M. Owada et al. / Precambrian Research 234 (2013) 63–84

Fig. 13. Chondrite-normalized REE patterns (a) and spider diagram (b) for the inferred primary melts of the low-Ti and high-Ti microgabbros. Normalized values of the chondrite and primitive mantle are quoted from Sun and McDonough (1989). These diagrams also show REE and trace elements abundances normalized to chondrite and primitive mantle. Note: The REE and trace element abundances for the inferred primary melts are similar to N-MORB (Sun and McDonough, 1989).

olivine. The above procedure added 8–13.5 wt% and 21.5–22 wt% olivines to the low-Ti and high-Ti microgabbros, respectively. Trace element compositions of primary melts are calculated from those of the estimated volume of olivine fractionation, using the formulation of Rayleigh fractionation equation. The chemical compositions of inferred primary magmas are listed in Table 8. Fig. 13 shows the chondrite normalized REE pattern and the primitive mantle normalized spider-diagram for the inferred primary melt compositions of the low-Ti and high-Ti microgabbros. The trace element abundances of the inferred primary melt of the low-Ti and high-Ti microgabbros exhibit depleted chemical compositions similar to N-MORB (Fig. 13a and b). The average ␧NdI values of low-Ti and high-Ti microgabbros are of +3.1 and +4.6, respectively. These geochemical signatures suggest that the primary melt of the low-Ti and high-Ti microgabbro could be derived from depleted source mantle. The source mantle of minette magma shows ␧NdI values that are −0.1 to +0.5 lower than those of inferred source mantles for the microgabbros (Fig. 8). Generally, a depleted mantle increases its ␧Nd value with time along the Nd evolution line; thereby, the

Fig. 14. Rb/Sr and Ce/Sm diagrams for inferred primary melts of the low-Ti and high-Ti microgabbros, the minette and the migmatite and pelitic gneiss.

␧Nd values of the source mantles that produced the low-Ti and high-Ti microgabbro would be higher than +3.0 at c. 560 Ma. The minette magma, therefore, has not been directly derived from the depleted mantle but from a more enriched mantle. In addition to high LILE/HFSE ratios, the minettes show convex shapes from Nd to Ti in the spider diagram (Fig. 5c and d). These patterns resemble those of the migmatites and the pelitic gneisses from the Sør Rondane Mountains (Fig. 5b and d). Fig. 14 plots Rb/Sr versus Ce/Sm for the inferred primary melt for the microgabbros, minettes, and migmatites and pelitic gneisses, indicating these rocks show a broadly positive correlation. In addition, the minettes are plotted between the inferred primary melt of the microgabbros and the migmatitic gneisses. The Rb/Sr and Ce/Sm ratios of basaltic magma close to primary compositions would reflect the source mantle because these elements indicate the high incompatibility against residual phases of the source mantle. Therefore, the source mantle of minette magma is supposed to be a mixture of depleted mantle and crustal materials. Considering the geochemical features, including trace elements and Nd isotopic ratios, it is concluded that the minette magma was produced by partial melting of a metasomatized mantle influenced by subducted crustal rocks such as the migmatite and the pelitic gneiss.

Fig. 15. Tectonic model of the formation of minette from a collisional event (a: 640–620 Ma) to extensional collapse and formation of the minette magma (b: c. 560 Ma). Crustal materials subducted at the level of upper mantle produced a metasomatized mantle (a). The minette magma was formed by the partial melting of the metasomatized mantle due to up welling of asthenosphere.

M. Owada et al. / Precambrian Research 234 (2013) 63–84

The minettes distinctly cut the foliations of high-grade gneisses and have chilled margins. Moreover, some minettes include vesicles or amygdales (Owada et al., 2010), suggesting that the minettes have intruded at a shallow level of the crust, probably close to the surface. The intrusive age of the minette (564 Ma) is the same as that of the Vengen granite (564 Ma; Shiraishi et al., 2008), which has geochemical features similar to the within-plate granite (Li et al., 2003). In addition, the Lunckeryggen granite, which is an intrusive rock coeval with the minette (Owada et al., 2010), is regarded as a post-collisional intrusion in terms of its geochemical signatures (Owada et al., 2006). Considering the occurrence, geochemistry, and geochronology of the minettes combined with the metamorphic ages, we can consider the timing of intrusion of minette magma as a post-cratonitized stage. The Latest Proterozoic magma activity (c. 560 Ma) in the Sør Rondane Mountains is characterized by bimodal magmatism, granites and minettes, indicating the external heat source derived from the mantle into the crust. Toyoshima et al. (1995) described the structural evolution of the central part of the Sør Rondane Mountains based on the characteristics of deformation, metamorphism and igneous activity. The granitic stocks and dikes intrude the final stage of the structural evolution, related to the extensional setting because some dikes show the fracture-filling signature. Moreover, the intruded host gneisses show an isothermal decompression path at this stage (Asami et al., 1992). This suggests that exhumation of the crust took place in the central part of the Sør Rondane Mountains during the latest Neoproterozoic magmatism that postdates the peak metamorphism by about 60–80 million years probably caused by a collisional event. This process could be explained by the generation of the magmas in an elevated geothermal gradient produced by extensional collapse, followed by the upwelling of asthenosphere and accompanied by mechanical thinning of the lithosphere. The tectonic model during this tectonothermal event is schematically shown in Fig. 15. Similar scenario was drawn in the late- to post-tectonic magmatism of the Central Dronning Maud Land as proposed by Jacobs et al. (2008), although the timing of magmatism in the Central Dronning Maud Land took place at the Early Cambrian (530 Ma), slightly later than that in the Sør Rondane Mountains (560 Ma). The geochemical studies, including Sr–Nd isotopic compositions, thus reveal that the microgabbros and minette originated from depleted and enriched sources, respectively. Consequently, the magmatic processes in the Sør Rondane Mountains reflect the evolution of the lithosphere during the Neoproterozoic—the depleted mantle at the initial subduction stage, subsequently evolving to an enriched mantle at the continental collision stage. This lithospheric evolution can, therefore, be explained by interaction between the depleted mantle and the enriched materials such as subducted crustal rocks.

Acknowledgments We wish to thank T. Hokada for many supports of field work and SHRIMP analysis. Thanks go to M. Abe and the members of 50th and 49th Japanese Antarctic Research Expedition for helpful discussion and advice during the field work. We also thank K. Shiraishi, Y. Motoyoshi, T. Kawasaki, Y. Hiroi, T. Sakiyama and D. Hyodo for valuable discussions. We acknowledge A. Hubert, G. Johnson-Amin, and member of the Belgian Antarctic Research Station (2008–2009). Thanks are due to Zilong Li and an anonymous reviewer for their critical reviews and many constructive comments, and also go to Satish-Kumar for useful comments and editorial handling. This work was supported by the Grant-in-Aid for Scientific Research provided by Japan Society for the Promotion of Science (23540559: M. Owada, 23540534: A. Kamei, 22244063: Y. Osanai).

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Appendix A. Supplementary data Supplementary data associated with this article can be found, in the online version, at http://dx.doi.org/10.1016/j.precamres. 2013.02.007.

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